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Revising Estimates of Aquatic Gross Oxygen Production by the Triple Oxygen Isotope Method to Incorporate… Manning, Cara C.; Howard, Evan M.; Nicholson, David P.; Ji, Brenda Y.; Sandwith, Zoe O.; Stanley, Rachel H. R. Oct 19, 2017

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Geophysical Research LettersRevising Estimates of Aquatic Gross Oxygen Productionby the Triple Oxygen Isotope Method to Incorporatethe Local Isotopic Composition of WaterCara C. Manning1,2,3,4 , EvanM. Howard1,2,3,5 , David P. Nicholson2 , Brenda Y. Ji6,Zoe O. Sandwith2, and Rachel H. R. Stanley61MIT/WHOI Joint Program in Oceanography/Applied Ocean Science and Engineering, Woods Hole, MA, USA, 2Departmentof Marine Chemistry and Geochemistry, Woods Hole Oceanographic Institution, Woods Hole, MA, USA, 3Department ofEarth, Atmospheric and Planetary Sciences, Massachusetts Institute of Technology, Cambridge, MA, USA, 4Department ofEarth, Ocean and Atmospheric Sciences, University of British Columbia, Vancouver, BC, Canada, 5School of Oceanography,University of Washington, Seattle, WA, USA, 6Department of Chemistry, Wellesley College, Wellesley, MA, USAAbstract Measurement of the triple oxygen isotope (TOI) composition of O2 is an established methodfor quantifying gross oxygen production (GOP) in natural waters. A standard assumption to this methodis that the isotopic composition of H2O, the substrate for photosynthetic O2, is equivalent to Vienna standardmean ocean water (VSMOW). We present and validate a method for estimating the TOI composition of H2Obased on mixing of local meteoric water and seawater H2O end-members, and incorporating the TOIcomposition of H2O into GOP estimates. In the ocean, GOP estimates based on assuming the H2O isequivalent to VSMOW can have systematic errors of up to 48% and in low-salinity systems, errors can bea factor of 2 or greater. In future TOI-based GOP studies, TOI measurements of O2 and H2O should bepaired when the H2O isotopic composition is expected to differ from VSMOW.1. IntroductionThe triple oxygen isotope (TOI) composition of O2 (i.e., the relative abundance of16O, 17O, and 18O) is a tracerof gross oxygen production (GOP) that has been used in marine environments globally over the past twodecades (Juranek &Quay, 2013; Luz & Barkan, 2000, 2005). GOP is the gross rate of photosynthetic O2 produc-tion by the ecosystem, prior to any respiratoryO2 consumption. Furthermore, the analyticalmethod generallyalso involves the measurement of O2/Ar gas ratios, which are used to quantify net oxygen production(NOP), that is, GOP minus community respiration from autotrophs and heterotrophs (Juranek & Quay, 2013).The approach enables simultaneous in situ estimation of GOP, NOP, and community respiration averaged overa time scale of days to weeks. The TOI approach exploits the unique TOI signatures of the two sources ofdissolved O2: photosynthetic O2, which is derived from H2O, and atmospheric O2, which is added to surfacewaters via air-water gas exchange and has a small mass-independent isotope signature (Barkan & Luz, 2011;Luz & Barkan, 2005). Although the method is O2-based, GOP and NOP can be converted from O2 to carbon(C) units (to gross primary production and net community production) using published stoichiometric ratios(Juranek & Quay, 2013; Laws, 1991).The TOI approach has some advantages over other techniques for simultaneously estimating oxygen produc-tion and respiration. For example, incubation-based techniques typically require assuming light respirationand dark respiration rates are equal and can be biased by “bottle effects,” and in situ measurements of dualoxygen isotopes (18O and 16O only) are sensitive to the choice of respiratory fractionation factor and cannotbe used in the open ocean as the rate of change in the tracers is too low (Hotchkiss et al., 2014; Staehr et al.,2012; Tobias et al., 2007).Since the TOI approach was first published in 1999, there have been frequent revisions to the calculationsto improve the accuracy of the GOP estimates (Kaiser, 2011; Luz and Barkan, 2005, 2011; Luz et al., 1999;Nicholson, 2011; Nicholson et al., 2014; Prokopenko et al., 2011). Here we demonstrate that the local relativeabundance of 16O, 17O, and 18O in H2O is an important variable affecting GOP calculations that is not typicallytaken into account in field-based studies of GOP. We present and validate a method for estimating the localRESEARCH LETTER10.1002/2017GL074375Key Points:• Systematic errors in gross oxygenproduction (GOP) estimates can occurdue to incorrect assumptions aboutthe water isotopic composition• We present a method for estimatingthe local oxygen isotopic compositionof water and incorporating it into GOPcalculations• Oxygen and water triple isotopemeasurements should be pairedin GOP studies when the isotopiccomposition of water differsfrom VSMOWSupporting Information:• Supporting Information S1Correspondence to:C. C. Manning,cmanning@whoi.eduCitation:Manning, C. C., Howard, E. M.,Nicholson, D. P., Ji, B. Y.,Sandwith, Z. O., & Stanley, R. H. R.(2017). Revising estimates of aquaticgross oxygen production by the tripleoxygen isotope method to incorporatethe local isotopic composition ofwater. Geophysical Research Letters, 44.https://doi.org/10.1002/2017GL074375Received 7 JUN 2017Accepted 22 SEP 2017Accepted article online 19 OCT 2017©2017. American Geophysical Union.All Rights Reserved.MANNING ET AL. GOP ESTIMATES INCORPORATING H2O ISOTOPES 1Geophysical Research Letters 10.1002/2017GL074375TOI composition of H2O (when paired measurements of O2 and H2O triple oxygen isotopes are not available)and incorporating the local H2O isotopic composition into GOP calculations. We demonstrate that thisapproach can be applied to accurately estimate GOP in environments where the isotopic composition of H2Odiffers from standardmean oceanwater. This approach is crucial to extending the application of the TOI tracerfrom the openocean,where it has been frequently used, to systems influencedbyprecipitation-derivedwaterwhich has a highly variable isotopic composition (Juranek & Quay, 2013; Jurikova et al., 2016; Luz et al., 2002;Yeung et al., 2012).2. CalculationsIn this section we provide background on TOI calculations and describe our method of estimating the localTOI compositionofH2Oand incorporating the TOI compositionofH2O into theGOPcalculation.MATLAB codeused for performing these calculations is available on GitHub (http://github.com/caramanning/calcGOP).The oxygen isotopic composition of a chemical species is characterized using𝛿18O = X18X18std− 1, (1)and expressed in ‰ by multiplying by 1,000. Here X18 is the measured 18O/16O ratio determined relative toX18std, the18O/16O ratioof a knownstandard. The term 𝛿17O is defined similarly, using 17O/16O ratios. Typically,O2in air is the standard for isotopic measurements of O2 (Juranek & Quay, 2013; Luz & Barkan, 2000). The currentbest practice is tonormalizeH2O isotopemeasurements using two standards: VSMOW(Vienna standardmeanocean water) and SLAP (standard light Antarctic precipitation) (Li et al., 2015; Schoenemann et al., 2013).On the VSMOW-SLAP scale, VSMOW has a 𝛿18O and 𝛿17O of 0‰. Hereafter, all O2 isotopic compositions arereferenced to air O2 and all H2O isotopic compositions are referenced to VSMOW-SLAP. In this paper, wedistinguish between the isotopic composition of the two substrates (O2 and H2O) as 𝛿18O-O2 and 𝛿18O-H2O.The typical assumption in GOP studies is that the water isotopic composition is equivalent to VSMOW, a stan-dard based on deep open ocean water (Barkan & Luz, 2011; Hamme et al., 2012; Nicholson, 2011). However,surface ocean water differs from VSMOWby up to 7‰ for 𝛿18O-H2O, though is most commonly within 2‰ofVSMOW (LeGrande & Schmidt, 2006). Furthermore, in brackish and freshwater systems the isotopic composi-tion of H2O can be substantially different from VSMOW (by up to ∼40‰ for 𝛿18O-H2O) because the isotopiccomposition ofmeteoric (precipitation derived) water has substantial regional and seasonal variability (Craig,1961; Epstein & Mayeda, 1953; Feng et al., 2009; Jasechko et al., 2013; Vachon et al., 2007).The term 17Δ quantifies the TOI composition of dissolved O2:17Δ = ln(𝛿17O-O2 + 1)− 𝜆 ln(𝛿18O-O2 + 1). (2)Here we report 17Δwith 𝜆 = 𝛾R, where𝛾R =17𝜖R18𝜖R= 0.5179. (3)In this equation, 𝜖R is the respiratory O2 consumption isotope effect (Luz & Barkan, 2005). When 𝜆 = 𝛾R,17Δreflects the proportion ofO2 that is derived fromphotosynthesis relative to air-water gas exchange (Hendrickset al., 2005; Juranek & Quay, 2013; Nicholson et al., 2014).We use 𝜆BSS to refer to the isotopic behavior of a system at biological steady state, in which the rates of photo-synthesis and respiration are approximately constant and within the same order of magnitude (Luz & Barkan,2005; Kaiser, 2011; Nicholson, 2011). The term 𝛾R describes the kinetic fractionation that occurs during respi-ratory O2 consumption and 𝜆BSS is an approximation of the trend in bulk O2 isotopic composition that arisesfrom respiratory fractionation and subsequent mixing between newly generated photosynthetic O2 and theresidual pool of O2 (e.g., see Figure 2 of Luz and Barkan (2005)). We calculate𝜆BSS =ln(17𝛼R)ln(18𝛼R) = ln(1 + 𝛾R18𝜖R)ln(1 + 18𝜖R) . (4)MANNING ET AL. GOP ESTIMATES INCORPORATING H2O ISOTOPES 2Geophysical Research Letters 10.1002/2017GL074375w = 0.528)biological steady state BSS = 0.5154)17BSS = 242ln(17 O + 1)H2OO2air O2photosynthetic O2VSMOWln( 18O + 1)17BSS > 24217BSS < 242 Figure 1. Schematic (not to scale), adapted from Luz and Barkan (2000),showing how the oxygen triple isotopic composition is modified as H2Ois converted to biological O2 in a steady state system with production,respiration, and no air-water gas exchange. The red line shows the oxygenisotopic fractionation as H2O with 𝛿18O-H2O = 0‰ and 17O-excess = 0 permeg versus VSMOW-SLAP (red circle) is converted to biological O2(red triangle) with 17Δ=242 per meg versus air O2. The light and darkblue lines show the oxygen isotopic fractionation of meteoric H2O, witha higher (dark blue) or lower (light blue) 𝛿18O-H2O than VSMOW, and a17O-excess of ∼30 per meg (Li et al., 2015; Luz & Barkan, 2000). The slope𝜆BSS is independent of the H2O isotopic composition whereas the valueof 17ΔBSS differs for each starting H2O isotopic composition.Assuming 18𝜖R=−20‰ (Kiddon et al., 1993) and 𝛾R=0.5179, 𝜆BSS=0.5154(Nicholson, 2011). Then we can calculate 17ΔBSS using 𝜆BSS17ΔBSS = ln(𝛿17O-O2 + 1)− 𝜆BSS ln(𝛿18O-O2 + 1). (5)Other authors have shown that 17ΔBSS is equivalent for pure photosyn-thetic O2 and for the bulk O2 isotopic composition after respiratory con-sumption and mixing (Angert et al., 2003; Nicholson, 2011).Using the isotopic compositionof photosyntheticO2 generatedby “averagephytoplankton” from VSMOW reported in Luz and Barkan (2011), the17Δ of biological O2 produced in a steady state system with production,respiration and no air-water gas exchange is ∼242 per meg (Figure 1).The 17Δ value for air-equilibrated water is ∼8 per meg (Reuer et al., 2007;Stanley et al., 2010). In natural waters, 17Δ of dissolved O2 will typicallyfall between these two end-members (Kaiser, 2011; Luz & Barkan, 2005;Nicholson, 2011).The term 17O-excess quantifies TOI composition of H2O:17O-excess = ln(𝛿17O-H2O + 1)− 𝜆W ln(𝛿18O-H2O + 1)(6)with𝜆W =0.528andall isotopic compositionsnormalized toVSMOW-SLAP.Equation (6) has a similar form to equation (5); however, the H2O isotopecommunity has traditionally used the 17O-excess terminology, whereasthe O2 isotope community has preferred the17Δ notation (Luz & Barkan,2010). The value of 𝜆W = 0.528 is well established for globally distributedmeasurements of meteoric waters and seawater (Landais et al., 2008; Luz& Barkan, 2010; Meijer & Li, 1998). Spatial variability in the 17O-excess ofnatural waters is less well understood relative to variability in 𝛿18O-O2 dueto the currently limited observations at sufficient accuracy to resolve the small excess (Barkan & Luz, 2005;Landais et al., 2008; Li et al., 2015; Luz & Barkan, 2010).In this paper, we quantify the effect of using the water isotopic composition in GOP calculations by com-bining measured 17Δ with previously published estimates of the local isotopic composition of H2O. We alsopresent a data set from the Arctic Ocean where 17Δ and 𝛿18O-H2O were measured simultaneously, and weestimate the 17O-excess. The approach used to estimate the isotopic composition of H2O for different loca-tions is based on the best available published data (LeGrande & Schmidt, 2006; Li et al., 2015; Luz & Barkan,2010; Schmidt et al., 1999). For brackish systems, we assume that the water sampled represents a mixtureof two end-members, seawater, and meteoric water, and use the sample salinity to estimate the fraction ofeach end-member (supporting information Text S1 and S2,Macdonald et al., 1995; Surge and Lohmann, 2002;Wankel et al., 2006).We calculate GOP using equation S8 from Prokopenko et al. (2011)GOP = kO2 [O2]eq[X17−X17eqX17− 𝜆X18−X18eqX18][X17P −X17X17− 𝜆 X18P −X18X18] + h[O2]𝜕17Δ𝜕t[X17P −X17X17− 𝜆 X18P −X18X18] . (7)Here kO2 is the gas transfer velocity for O2 (m d−1), [O2] is the O2 concentration (mol m−3), h is the mixedlayer depth (m), X17 is the 17O∕16O ratio of O2, and the subscripts eq and P refer to O2 at air-water equilibriumand O2 produced by photosynthesis, respectively. Oxygen terms without a subscript (O2, X17, and 17Δ) arethe measured mixed layer values. We calculate X18eq based on Benson & Krause (1980, 1984) and then deriveX17eq using17Δeq = 8 per meg (Reuer et al., 2007; Stanley et al., 2010). We calculate X18P and X17P using the iso-topic composition of H2O estimated as described above, and isotopic fractionation factors for photosyntheticO2 with respect to H2O based on the average of four different types of phytoplankton, using data in Table 1MANNING ET AL. GOP ESTIMATES INCORPORATING H2O ISOTOPES 3Geophysical Research Letters 10.1002/2017GL074375of Luz and Barkan (2011), specifically 18𝜖P = 3.389‰ and 17𝜖P = 1.778‰. We use 𝜆 = 𝛾R = 0.5179 inequation (7) (supporting information Text S3, Hamme et al., 2012; Juranek and Quay, 2013; Kaiser, 2011;Nicholson, 2011; Nicholson et al., 2012). The triple isotopic composition of photosynthetic O2 can be easilymodified in equation (7) and our software for calculating GOP. Such modifications could be required due toongoing revisions in the isotopic composition of atmospheric O2 with respect to VSMOW (Barkan & Luz, 2011;Pack et al., 2016; Young et al., 2014) or changes in the photosynthetic fractionation factors to reflect a differ-ent community composition (section 3) and/or new experimental data (Angert et al., 2003; Eisenstadt et al.,2010; Luz & Barkan, 2011).The first term on the right side of equation (7) is the steady state GOP term, and the second term is the non-steady stateGOP term. If there is no change in 17Δwith time (steady statewith respect to 17Δ), then the secondterm on the right side of equation (7) equals zero. The relative (percent) error in GOP associated with the localH2O isotopic composition is independent of the magnitude of the nonsteady state term. Additional terms tocorrect for physical transport andmixing processes can also be included in the GOP calculation (Munro et al.,2013; Nicholson et al., 2014;Wurgaft et al., 2013), but we omit further discussion of these corrections here andinstead focus on the sensitivity of the GOP calculations to the H2O isotopic composition.3. Chamber ExperimentsWe conducted field-based incubation experiments which demonstrate that the TOI composition of O2 in asystem approximating biological steady state agreedwith the value predicted from the local isotopic compo-sition of water, rather than the typically used VSMOW value. These experiments used benthic flux chambers(Stanley&Howard, 2013) in tidal creeks in thePlum IslandEstuary,Massachusetts. The chamberswere insertedinto thewater-sediment interface and sealed to limit exchangewith surrounding creekwater (Howard, 2017).Over the ∼4 h daytime incubation period, the concentration of O2 in each chamber increased due to photo-synthetic production exceeding respiration, and the fractionofO2 derived fromphotosynthesis also increased(Figure 2). Based on the initial O2 isotopic composition, observedO2 concentrations in the chambers, and esti-mates of ecosystem production and respiration, it is likely that by the end of the incubation >90% of the O2in each chamber was biological in origin (Howard, 2017), and therefore 17Δ of the dissolved O2 approachedthe biological steady state value (Luz & Barkan, 2000). By the end of each incubation, the value of 17Δ(calculated with 𝜆=0.5179) stabilized at a value near 210 per meg, rather than continuing to increase towardthe value of ∼242 per meg predicted for a steady state system where photosynthetic O2 is derived fromVSMOW (Figure 2).Using an estimated local isotopic composition of H2O based on published data and the salinity in the cham-bers (supporting information Text S1), we predict 17ΔBSS (calculated with 𝜆BSS = 0.5154) to be 215 per megassuming the isotopic fractionation factor for “average phytoplankton” (average of diatoms, cyanobacteria,coccolithophores, and green algae) defined in Luz and Barkan (2011) or 206 per meg assuming a 1:1 mixtureof diatoms and cyanobacteria only (based on microbial cultures from water sampled in the creeks). Thesepredicted values are consistent with the observed 17Δ of ∼210 per meg.If the incubations were extended for a longer duration, all of the atmospheric O2 should eventually beremoved by respiration and the O2 within the chamber should become essentially 100% biological in originand have a composition equivalent to 17ΔBSS (206–215 permeg) (Luz & Barkan, 2011; Nicholson, 2011). If 90%of the O2 in the chamber were biological in origin at the end of the incubation (and the rest derived from gasexchange), the expected 17ΔBSS would be 194–205 per meg.Because the measured 17ΔBSS matched the predicted value with the estimated local isotopic composition ofH2O taken into account, we conclude that chamber incubations are an effective approach for determining17ΔBSS in coastal waters. There are several advantages to the chamber approach compared to published tech-niques for estimating 17ΔBSS (Luz & Barkan, 2011; Luz et al., 1999). For example, the method measures theisotopic composition of dissolvedO2 directly, rather than the composition of gaseous O2 in a headspace equi-librated with water (Luz et al., 1999; Luz & Barkan, 2000), the phytoplankton are grown in oxygenated water(not zero-O2 conditions) (Barkan & Luz, 2011; Eisenstadt et al., 2010; Luz & Barkan, 2011), and we determinethe isotopic composition for an intact microbial community in situ rather than a single autotrophic species inlab-based culture.MANNING ET AL. GOP ESTIMATES INCORPORATING H2O ISOTOPES 4Geophysical Research Letters 10.1002/2017GL074375Figure 2. (a) Observed and predicted measurements of 17Δ during three chamber incubation experiments. The coloredlines show measured values of 17Δ (calculated with 𝜆 = 𝛾R = 0.5179). The gray shaded region shows the predictedvalue of 17ΔBSS (calculated with 𝜆 = 𝜆BSS = 0.5154) using the estimated local H2O isotopic composition, and the dashedblack line shows the 17ΔBSS value for VSMOW water (see text). Error bars indicate uncertainty due to measurementand sampling. (b) Measured O2 concentrations in each chamber. Dash-dotted black line shows the equilibrium O2concentration. Variability in [O2] between chambers is due to expected heterogeneity in benthic estuarine sediments.Notably, 𝛿18O-O2 and 𝛿17O-O2 approached an asymptote in the same manner as17Δ but differed betweenchambers and did not reach the published values for photosynthetic end-members (supporting informa-tion Text S4 and Figure S1). GOP calculations are much more sensitive to 17ΔBSS than to the absolute valueof 𝛿18O-O2P . For example, changing 𝛿18O-O2P from −22‰ to −11‰ changes GOP by less than 5%, as longas 17ΔBSS is held constant. Therefore, the fact that we estimated 17ΔBSS but not 𝛿18O-O2P does not signifi-cantly reduce the utility of the chamber incubation approach and does not affect our conclusion that thelocal TOI composition of H2O should be incorporated intoGOP calculations in systemswhere theH2O isotopiccomposition differs from VSMOW.4. Beaufort Sea, Arctic OceanIn the Arctic Ocean, 𝛿18O-H2O in the surface mixed layer can range from −7 to 0‰ (LeGrande & Schmidt,2006). To quantify the effect of incorporating the local H2O isotopic composition into Arctic Ocean GOP esti-mates, we use a data set of measurements of 17Δ from summer 2011 to 2012 in the Beaufort Sea region ofthe Canada Basin (Stanley et al., 2015) and incorporate measurements of 𝛿18O-H2O from the same cruise. Weassumeafixed 17O-excess of−5permeg, the averagenear-surface seawater value fromLuz andBarkan (2010);however, this value is uncertain because no published observations are available in this region. For this dataset, the calculated GOP assuming the isotopic composition of H2O is VSMOW is on average 29% lower (range48 to 19% lower) than the calculated GOP using the local isotopic composition of H2O (Figures 3 and S2 andTable 1). Increasing or decreasing the 17O-excess by 8 per meg for all samples (two standard deviations ofthe global seawater data set from Luz and Barkan (2010)) changes the calculated GOP by 5–6% on average.The systematic error in Beaufort Sea GOP calculations associated with assuming the H2O isotopic composi-tion is equivalent to VSMOW is larger in magnitude than most other sources of error in the calculation, suchas the gas transfer velocity (10–15% error) (Ho et al., 2006). Given that the isotopic composition of H2O dif-fers from VSMOW throughoutmost of the Arctic Ocean, we expect that GOP estimates in other regions of theArctic Ocean would have similarly large systematic errors if the isotopic composition of H2O is not incorpo-rated into the GOP calculation.For future studies of GOP from 17Δmeasurements in the Arctic Ocean, we strongly recommend investigatorsalsomeasure 𝛿18O-H2O and preferably the17O-excess and that they include these variables in the GOP calcu-lation. Recently, in many regions of the Arctic Ocean the freshwater content has substantially increased andsea ice cover has substantially decreased, and therefore, it would likely be inaccurate to infer the TOI com-position of H2O using previously published, location-based measurements of 𝛿18O-H2O (McPhee et al., 2009;Morison et al., 2012; Yamamoto-Kawai et al., 2009). Additionally, in the Arctic Ocean, there are three salin-ity and H2O isotope end-members (seawater, meteoric water, and sea ice melt) (Eicken et al., 2002; Ekwurzelet al., 2001; Östlund & Hut, 1984; Yamamoto-Kawai et al., 2008). The 17O-excess of these three end-membersMANNING ET AL. GOP ESTIMATES INCORPORATING H2O ISOTOPES 5Geophysical Research Letters 10.1002/2017GL074375-8 -6 -4 -2 0 2δ18O-H2O [per mil vs. VSMOW-SLAP]-60-50-40-30-20-10010% error in GOPVSMOW0 50 100 15017 [per meg vs. air O2]Beaufort SeaBering SeaNorth PacificBellingshausenCaliforniaSouth AtlanticWaquoit Bay-70-60-60-50-40-40-30-30-20-20-10-1000101020-6 -5 -4 -3 -2 -1 0 1 2δ18O-H2O [per mil vs. VSMOW-SLAP]02040608010012014017 [per meg vs. air O2]Beaufort SeaBering SeaNorth PacificBellingshausenCaliforniaSouth Atlantic-60-40-20020% error inGOPVSMOWa           bc-50Figure 3. Error in GOP calculations caused by assuming the water isotopic composition is equivalent to VSMOW(GOPVSMOW) compared to using the local H2O isotopic composition. Percent error as a function of (a) 𝛿18O-H2O and(b) 17Δ for all data sets shown in Table 1. Black lines indicate one standard deviation of the values for each region.(c) Contour plot showing % error in GOPVSMOW as a function of17Δ and 𝛿18O-H2O, for a fixed 17O-excess of −5 per meg(the average oceanic value), calculated using equation (7). Individual samples from each oceanic data set are shown.Table 1Globally Distributed Triple Oxygen Isotope Data and Errors in Calculated GOP Caused by Assuming the H2O Isotopic Composition is Equivalent to VSMOW𝛿18O-H2Oa 17O-excessa(‰ versus (per meg versus 17ΔLocation VSMOW-SLAP) VSMOW-SLAP) (per meg versus air O2) % Error in GOPb ReferenceBeaufort Sea −3.6 (−6.1, −2.3) −5 22 (6, 80) −29 (−48, −19) Stanley et al. (2015)Bering Sea −1.1 (−1.7, −0.7) −5 45 (8, 119) −11 (−16, −7) Prokopenko et al. (2011)California Coast −0.5 (−0.5, −0.3) −5 64 (14, 170) −6 (−11, −5) Munro et al. (2013)Northeast Pacific −0.4 (−0.8, 0.4) −5 34 (11, 67) −4 (−9, 1) Juranek et al. (2012)South Atlantic 0.8 (0.3, 1.4) −5 32 (17, 50) 5 (1, 6) Howard (2017)Bellingshausen Sea −0.5 (−0.7, −0.4) −5 35 (9, 109) −6 (−10, −5) Castro-Morales et al. (2013)Bras d’Or Lake −7.1 (−9.3, −4.8) 16 (6, 27) 35 (9, 54) −46 (−56, −33) Manning (2017)Waquoit Bay −2.9 (−6.1, −1.1) 1 (−5, 11) 100 (45, 236) −30 (−234, −11) This manuscriptNotes. Values are reported as mean (min, max). All tabulated data include only near-surface samples (depth <10 m). aDetails on the estimation of 𝛿18O-H2O and17O-excess for eachdata set are provided in the supporting information Text S1. Analyticalmethods are described in supporting information Text S5 and S6 (Barkanand Luz, 2003; Epstein and Mayeda, 1953; Stanley et al., 2010, 2015). bPercent error in GOP is calculated as % error = (GOPVSMOW − GOP)∕GOP × 100% whereGOPVSMOW is calculated assuming the H2O is equivalent to VSMOW and GOP is calculated using the estimated local H2O isotopic composition.MANNING ET AL. GOP ESTIMATES INCORPORATING H2O ISOTOPES 6Geophysical Research Letters 10.1002/2017GL074375is currently unknown in this region. Therefore, it is not possible to use the two-end-member approach andestimate 𝛿18O-H2O and 𝛿17O-H2O based on salinity alone, nor to estimate 𝛿17O-H2O based on measured𝛿18O-H2O and salinity (supporting information Text S1). If the17O-excess of the three end-members wereknown, one could estimate the 17O-excess using a water mass analysis based on measurements of 𝛿18O-H2Oand an additional freshwater source tracer such as total alkalinity (Yamamoto-Kawai et al., 2008). Concurrentmeasurements of both 𝛿18O-H2O and17O-excess would avoid the uncertainties associated with the watermass analysis.5. Implications for Oceanic and Estuarine SystemsWe use a variety of TOI measurements in other oceanic regions and brackish waters to estimate the error inGOP estimates caused by assuming the water is equivalent to VSMOW (Table 1). In all data sets, the magni-tude of the error in GOP estimates is strongly dependent on the magnitude of 𝛿18O-H2O and more weaklydependent on the magnitude of 17Δ (Figure 3c). In the Bering Sea, GOP is up to 16% lower when calculatedusing VSMOW H2O instead of the local H2O isotopic composition, and in other oceanic regions outside theArctic, the maximum error varies from 11% too low to 6% too high.In freshwater and brackish regions, the magnitude of the errors in calculated GOP is significantly larger, up to56% in an estuary inNova Scotia, Canada, andover 200% in an estuary inMassachusetts (Table 1). These largererrors occur primarily due to the large influence of meteoric water in estuaries, which has a 𝛿18O-H2O of −6to −9‰ in these regions, and in the Massachusetts estuary the errors are exacerbated by high values of 17Δin some samples (Figure 3b). Notably, there are somewhat large uncertainties in the appropriate 17O-excessto use for brackish systems because published data are limited and the range of observed values is large(supporting information Text S2). In future studies, we recommend that authors applying the TOI approachto estimate GOP in regions influenced by meteoric water should measure the 17O-excess of H2O in additionto 17Δ of O2.6. Conclusions and Future OutlookWe have shown that the local TOI composition of H2O is an important variable to include in GOP calculationsin systems where the isotopic composition of H2O differs from VSMOW. We have presented a quantita-tive approach for estimating the local TOI composition of H2O and incorporating this variable into GOPcalculations. Chamber incubation data in an estuary validated this approach and suggest that the TOIapproach could be used to estimate GOP in estuarine and freshwater systems, which could greatly expand itsutility (Staehr et al., 2012).New analytical methods could enable the routine measurement of 𝛿18O-H2O and17O-excess as a part offuture TOI-based GOP studies and promise to increase scientific understanding of spatial trends in H2Oisotopes. Until the past couple of years, analytical methods with sufficient precision to resolve the 17O-excessof H2O were mass spectrometer-based, required a specialized front end and the generation of hazardouschemicals, and did not permit the simultaneous measurement of 𝛿17O and 𝛿2H (Luz & Barkan, 2005).The recent development of commercially available laser-based sensors that simultaneously measure the17O-excess and 𝛿2H of H2O directly, without complicated sample processing (Berman et al., 2013; Steig et al.,2014), will greatly increase the number of laboratories making this measurement and reduce analytical costs(Schauer et al., 2016).Openly accessible databases of 17O-excess in natural waters, similar to existing databases for 𝛿18O-H2O and𝛿2H-H2O (Schmidt et al., 1999; IAEA/WMO, 2016), would be highly beneficial to the stable isotope geochem-istry community. For example, these data could improve estimates of the 17O-excess for data sets where TOImeasurements of O2 and H2O were not collected simultaneously, including previously published data sets.Additionally, public archiving of all data used in TOI-based GOP studies will ensure that GOP can be recalcu-lated in the future based on new consensus about the most appropriate way to perform these calculations(supporting information Text S2).ReferencesAngert, A., Rachmilevitch, S., Barkan, E., & Luz, B. (2003). Effects of photorespiration, the cytochrome pathway, and the alternative pathwayon the triple isotopic composition of atmospheric O2. Global Biogeochemical Cycles, 17(1), 1030. https://doi.org/10.1029/2002GB001933Barkan, E., & Luz, B. (2003). High-precision measurements of 17O/16O and 18O/16O of O2 and O2/Ar ratio in air. Rapid Communications inMass Spectrometry, 17(24), 2809–2814.AcknowledgmentsWe acknowledge funding fromNSF OCE-1029676, OCE-1233678,PLR-1304406, and PLR-1547011 toR.H.R. Stanley, scholarships fromNSERC and CMOS to C.C. Manning,and an NDSEG fellowship toE. M. Howard. We thank MichiyoYamamoto-Kawai for providing theCanada Basin 𝛿18O-H2O data andAmanda Spivak for assisting withincubation chamber design. We thankBoaz Luz and Laurence Yeung for theirthoughtful reviews that improved themanuscript. All data sets collectedby the coauthors that were used inthis study are freely available online:https://doi.org/10.18739/A2W389for Beaufort Sea, https://doi.org/10.5281/zenodo.351764 for WaquoitBay, https://doi.org/10.5281/zenodo.352786 for Bras d’Or Lake, andhttps://doi.org/10.1575/1912/8654 forPlum Island Estuary and the SouthAtlantic. Software for calculatingGOP is available on GitHub(Manning and Howard (2017),http://github.com/caramanning/calcGOP).MANNING ET AL. 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D., Yeung, L. Y., & Kohl, I. E. (2014). On theΔ17O budget of atmospheric O2. Geochimica et Cosmochimica Acta, 135, 102–125.MANNING ET AL. GOP ESTIMATES INCORPORATING H2O ISOTOPES 9 Geophysical Research LettersRevising Estimates of Aquatic Gross Oxygen Productionby the Triple Oxygen Isotope Method to Incorporatethe Local Isotopic Composition of WaterCara C. Manning1,2,3,4 , EvanM. Howard1,2,3,5 , David P. Nicholson2 , Brenda Y. Ji6,Zoe O. Sandwith2, and Rachel H. R. Stanley61MIT/WHOI Joint Program in Oceanography/Applied Ocean Science and Engineering, Woods Hole, MA, USA, 2Departmentof Marine Chemistry and Geochemistry, Woods Hole Oceanographic Institution, Woods Hole, MA, USA, 3Department ofEarth, Atmospheric and Planetary Sciences, Massachusetts Institute of Technology, Cambridge, MA, USA, 4Department ofEarth, Ocean and Atmospheric Sciences, University of British Columbia, Vancouver, BC, Canada, 5School of Oceanography,University of Washington, Seattle, WA, USA, 6Department of Chemistry, Wellesley College, Wellesley, MA, USAAbstract Measurement of the triple oxygen isotope (TOI) composition of O2 is an established methodfor quantifying gross oxygen production (GOP) in natural waters. A standard assumption to this methodis that the isotopic composition of H2O, the substrate for photosynthetic O2, is equivalent to Vienna standardmean ocean water (VSMOW). We present and validate a method for estimating the TOI composition of H2Obased on mixing of local meteoric water and seawater H2O end-members, and incorporating the TOIcomposition of H2O into GOP estimates. In the ocean, GOP estimates based on assuming the H2O isequivalent to VSMOW can have systematic errors of up to 48% and in low-salinity systems, errors can bea factor of 2 or greater. In future TOI-based GOP studies, TOI measurements of O2 and H2O should bepaired when the H2O isotopic composition is expected to differ from VSMOW.1. IntroductionThe triple oxygen isotope (TOI) composition of O2 (i.e., the relative abundance of16O, 17O, and 18O) is a tracerof gross oxygen production (GOP) that has been used in marine environments globally over the past twodecades (Juranek &Quay, 2013; Luz & Barkan, 2000, 2005). GOP is the gross rate of photosynthetic O2 produc-tion by the ecosystem, prior to any respiratoryO2 consumption. Furthermore, the analyticalmethod generallyalso involves the measurement of O2/Ar gas ratios, which are used to quantify net oxygen production(NOP), that is, GOP minus community respiration from autotrophs and heterotrophs (Juranek & Quay, 2013).The approach enables simultaneous in situ estimation of GOP, NOP, and community respiration averaged overa time scale of days to weeks. The TOI approach exploits the unique TOI signatures of the two sources ofdissolved O2: photosynthetic O2, which is derived from H2O, and atmospheric O2, which is added to surfacewaters via air-water gas exchange and has a small mass-independent isotope signature (Barkan & Luz, 2011;Luz & Barkan, 2005). Although the method is O2-based, GOP and NOP can be converted from O2 to carbon(C) units (to gross primary production and net community production) using published stoichiometric ratios(Juranek & Quay, 2013; Laws, 1991).The TOI approach has some advantages over other techniques for simultaneously estimating oxygen produc-tion and respiration. For example, incubation-based techniques typically require assuming light respirationand dark respiration rates are equal and can be biased by “bottle effects,” and in situ measurements of dualoxygen isotopes (18O and 16O only) are sensitive to the choice of respiratory fractionation factor and cannotbe used in the open ocean as the rate of change in the tracers is too low (Hotchkiss et al., 2014; Staehr et al.,2012; Tobias et al., 2007).Since the TOI approach was first published in 1999, there have been frequent revisions to the calculationsto improve the accuracy of the GOP estimates (Kaiser, 2011; Luz and Barkan, 2005, 2011; Luz et al., 1999;Nicholson, 2011; Nicholson et al., 2014; Prokopenko et al., 2011). Here we demonstrate that the local relativeabundance of 16O, 17O, and 18O in H2O is an important variable affecting GOP calculations that is not typicallytaken into account in field-based studies of GOP. We present and validate a method for estimating the localRESEARCH LETTER10.1002/2017GL074375Key Points:• Systematic errors in gross oxygenproduction (GOP) estimates can occurdue to incorrect assumptions aboutthe water isotopic composition• We present a method for estimatingthe local oxygen isotopic compositionof water and incorporating it into GOPcalculations• Oxygen and water triple isotopemeasurements should be pairedin GOP studies when the isotopiccomposition of water differsfrom VSMOWSupporting Information:• Supporting Information S1Correspondence to:C. C. Manning,cmanning@whoi.eduCitation:Manning, C. C., Howard, E. M.,Nicholson, D. P., Ji, B. Y.,Sandwith, Z. O., & Stanley, R. H. R.(2017). Revising estimates of aquaticgross oxygen production by the tripleoxygen isotope method to incorporatethe local isotopic composition ofwater. Geophysical Research Letters, 44.https://doi.org/10.1002/2017GL074375Received 7 JUN 2017Accepted 22 SEP 2017Accepted article online 19 OCT 2017©2017. American Geophysical Union.All Rights Reserved.MANNING ET AL. GOP ESTIMATES INCORPORATING H2O ISOTOPES 1Geophysical Research Letters 10.1002/2017GL074375TOI composition of H2O (when paired measurements of O2 and H2O triple oxygen isotopes are not available)and incorporating the local H2O isotopic composition into GOP calculations. We demonstrate that thisapproach can be applied to accurately estimate GOP in environments where the isotopic composition of H2Odiffers from standardmean oceanwater. This approach is crucial to extending the application of the TOI tracerfrom the openocean,where it has been frequently used, to systems influencedbyprecipitation-derivedwaterwhich has a highly variable isotopic composition (Juranek & Quay, 2013; Jurikova et al., 2016; Luz et al., 2002;Yeung et al., 2012).2. CalculationsIn this section we provide background on TOI calculations and describe our method of estimating the localTOI compositionofH2Oand incorporating the TOI compositionofH2O into theGOPcalculation.MATLAB codeused for performing these calculations is available on GitHub (http://github.com/caramanning/calcGOP).The oxygen isotopic composition of a chemical species is characterized using𝛿18O = X18X18std− 1, (1)and expressed in ‰ by multiplying by 1,000. Here X18 is the measured 18O/16O ratio determined relative toX18std, the18O/16O ratioof a knownstandard. The term 𝛿17O is defined similarly, using 17O/16O ratios. Typically,O2in air is the standard for isotopic measurements of O2 (Juranek & Quay, 2013; Luz & Barkan, 2000). The currentbest practice is tonormalizeH2O isotopemeasurements using two standards: VSMOW(Vienna standardmeanocean water) and SLAP (standard light Antarctic precipitation) (Li et al., 2015; Schoenemann et al., 2013).On the VSMOW-SLAP scale, VSMOW has a 𝛿18O and 𝛿17O of 0‰. Hereafter, all O2 isotopic compositions arereferenced to air O2 and all H2O isotopic compositions are referenced to VSMOW-SLAP. In this paper, wedistinguish between the isotopic composition of the two substrates (O2 and H2O) as 𝛿18O-O2 and 𝛿18O-H2O.The typical assumption in GOP studies is that the water isotopic composition is equivalent to VSMOW, a stan-dard based on deep open ocean water (Barkan & Luz, 2011; Hamme et al., 2012; Nicholson, 2011). However,surface ocean water differs from VSMOWby up to 7‰ for 𝛿18O-H2O, though is most commonly within 2‰ofVSMOW (LeGrande & Schmidt, 2006). Furthermore, in brackish and freshwater systems the isotopic composi-tion of H2O can be substantially different from VSMOW (by up to ∼40‰ for 𝛿18O-H2O) because the isotopiccomposition ofmeteoric (precipitation derived) water has substantial regional and seasonal variability (Craig,1961; Epstein & Mayeda, 1953; Feng et al., 2009; Jasechko et al., 2013; Vachon et al., 2007).The term 17Δ quantifies the TOI composition of dissolved O2:17Δ = ln(𝛿17O-O2 + 1)− 𝜆 ln(𝛿18O-O2 + 1). (2)Here we report 17Δwith 𝜆 = 𝛾R, where𝛾R =17𝜖R18𝜖R= 0.5179. (3)In this equation, 𝜖R is the respiratory O2 consumption isotope effect (Luz & Barkan, 2005). When 𝜆 = 𝛾R,17Δreflects the proportion ofO2 that is derived fromphotosynthesis relative to air-water gas exchange (Hendrickset al., 2005; Juranek & Quay, 2013; Nicholson et al., 2014).We use 𝜆BSS to refer to the isotopic behavior of a system at biological steady state, in which the rates of photo-synthesis and respiration are approximately constant and within the same order of magnitude (Luz & Barkan,2005; Kaiser, 2011; Nicholson, 2011). The term 𝛾R describes the kinetic fractionation that occurs during respi-ratory O2 consumption and 𝜆BSS is an approximation of the trend in bulk O2 isotopic composition that arisesfrom respiratory fractionation and subsequent mixing between newly generated photosynthetic O2 and theresidual pool of O2 (e.g., see Figure 2 of Luz and Barkan (2005)). We calculate𝜆BSS =ln(17𝛼R)ln(18𝛼R) = ln(1 + 𝛾R18𝜖R)ln(1 + 18𝜖R) . (4)MANNING ET AL. GOP ESTIMATES INCORPORATING H2O ISOTOPES 2Geophysical Research Letters 10.1002/2017GL074375w = 0.528)biological steady state BSS = 0.5154)17BSS = 242ln(17 O + 1)H2OO2air O2photosynthetic O2VSMOWln( 18O + 1)17BSS > 24217BSS < 242 Figure 1. Schematic (not to scale), adapted from Luz and Barkan (2000),showing how the oxygen triple isotopic composition is modified as H2Ois converted to biological O2 in a steady state system with production,respiration, and no air-water gas exchange. The red line shows the oxygenisotopic fractionation as H2O with 𝛿18O-H2O = 0‰ and 17O-excess = 0 permeg versus VSMOW-SLAP (red circle) is converted to biological O2(red triangle) with 17Δ=242 per meg versus air O2. The light and darkblue lines show the oxygen isotopic fractionation of meteoric H2O, witha higher (dark blue) or lower (light blue) 𝛿18O-H2O than VSMOW, and a17O-excess of ∼30 per meg (Li et al., 2015; Luz & Barkan, 2000). The slope𝜆BSS is independent of the H2O isotopic composition whereas the valueof 17ΔBSS differs for each starting H2O isotopic composition.Assuming 18𝜖R=−20‰ (Kiddon et al., 1993) and 𝛾R=0.5179, 𝜆BSS=0.5154(Nicholson, 2011). Then we can calculate 17ΔBSS using 𝜆BSS17ΔBSS = ln(𝛿17O-O2 + 1)− 𝜆BSS ln(𝛿18O-O2 + 1). (5)Other authors have shown that 17ΔBSS is equivalent for pure photosyn-thetic O2 and for the bulk O2 isotopic composition after respiratory con-sumption and mixing (Angert et al., 2003; Nicholson, 2011).Using the isotopic compositionof photosyntheticO2 generatedby “averagephytoplankton” from VSMOW reported in Luz and Barkan (2011), the17Δ of biological O2 produced in a steady state system with production,respiration and no air-water gas exchange is ∼242 per meg (Figure 1).The 17Δ value for air-equilibrated water is ∼8 per meg (Reuer et al., 2007;Stanley et al., 2010). In natural waters, 17Δ of dissolved O2 will typicallyfall between these two end-members (Kaiser, 2011; Luz & Barkan, 2005;Nicholson, 2011).The term 17O-excess quantifies TOI composition of H2O:17O-excess = ln(𝛿17O-H2O + 1)− 𝜆W ln(𝛿18O-H2O + 1)(6)with𝜆W =0.528andall isotopic compositionsnormalized toVSMOW-SLAP.Equation (6) has a similar form to equation (5); however, the H2O isotopecommunity has traditionally used the 17O-excess terminology, whereasthe O2 isotope community has preferred the17Δ notation (Luz & Barkan,2010). The value of 𝜆W = 0.528 is well established for globally distributedmeasurements of meteoric waters and seawater (Landais et al., 2008; Luz& Barkan, 2010; Meijer & Li, 1998). Spatial variability in the 17O-excess ofnatural waters is less well understood relative to variability in 𝛿18O-O2 dueto the currently limited observations at sufficient accuracy to resolve the small excess (Barkan & Luz, 2005;Landais et al., 2008; Li et al., 2015; Luz & Barkan, 2010).In this paper, we quantify the effect of using the water isotopic composition in GOP calculations by com-bining measured 17Δ with previously published estimates of the local isotopic composition of H2O. We alsopresent a data set from the Arctic Ocean where 17Δ and 𝛿18O-H2O were measured simultaneously, and weestimate the 17O-excess. The approach used to estimate the isotopic composition of H2O for different loca-tions is based on the best available published data (LeGrande & Schmidt, 2006; Li et al., 2015; Luz & Barkan,2010; Schmidt et al., 1999). For brackish systems, we assume that the water sampled represents a mixtureof two end-members, seawater, and meteoric water, and use the sample salinity to estimate the fraction ofeach end-member (supporting information Text S1 and S2,Macdonald et al., 1995; Surge and Lohmann, 2002;Wankel et al., 2006).We calculate GOP using equation S8 from Prokopenko et al. (2011)GOP = kO2 [O2]eq[X17−X17eqX17− 𝜆X18−X18eqX18][X17P −X17X17− 𝜆 X18P −X18X18] + h[O2]𝜕17Δ𝜕t[X17P −X17X17− 𝜆 X18P −X18X18] . (7)Here kO2 is the gas transfer velocity for O2 (m d−1), [O2] is the O2 concentration (mol m−3), h is the mixedlayer depth (m), X17 is the 17O∕16O ratio of O2, and the subscripts eq and P refer to O2 at air-water equilibriumand O2 produced by photosynthesis, respectively. Oxygen terms without a subscript (O2, X17, and 17Δ) arethe measured mixed layer values. We calculate X18eq based on Benson & Krause (1980, 1984) and then deriveX17eq using17Δeq = 8 per meg (Reuer et al., 2007; Stanley et al., 2010). We calculate X18P and X17P using the iso-topic composition of H2O estimated as described above, and isotopic fractionation factors for photosyntheticO2 with respect to H2O based on the average of four different types of phytoplankton, using data in Table 1MANNING ET AL. GOP ESTIMATES INCORPORATING H2O ISOTOPES 3Geophysical Research Letters 10.1002/2017GL074375of Luz and Barkan (2011), specifically 18𝜖P = 3.389‰ and 17𝜖P = 1.778‰. We use 𝜆 = 𝛾R = 0.5179 inequation (7) (supporting information Text S3, Hamme et al., 2012; Juranek and Quay, 2013; Kaiser, 2011;Nicholson, 2011; Nicholson et al., 2012). The triple isotopic composition of photosynthetic O2 can be easilymodified in equation (7) and our software for calculating GOP. Such modifications could be required due toongoing revisions in the isotopic composition of atmospheric O2 with respect to VSMOW (Barkan & Luz, 2011;Pack et al., 2016; Young et al., 2014) or changes in the photosynthetic fractionation factors to reflect a differ-ent community composition (section 3) and/or new experimental data (Angert et al., 2003; Eisenstadt et al.,2010; Luz & Barkan, 2011).The first term on the right side of equation (7) is the steady state GOP term, and the second term is the non-steady stateGOP term. If there is no change in 17Δwith time (steady statewith respect to 17Δ), then the secondterm on the right side of equation (7) equals zero. The relative (percent) error in GOP associated with the localH2O isotopic composition is independent of the magnitude of the nonsteady state term. Additional terms tocorrect for physical transport andmixing processes can also be included in the GOP calculation (Munro et al.,2013; Nicholson et al., 2014;Wurgaft et al., 2013), but we omit further discussion of these corrections here andinstead focus on the sensitivity of the GOP calculations to the H2O isotopic composition.3. Chamber ExperimentsWe conducted field-based incubation experiments which demonstrate that the TOI composition of O2 in asystem approximating biological steady state agreedwith the value predicted from the local isotopic compo-sition of water, rather than the typically used VSMOW value. These experiments used benthic flux chambers(Stanley&Howard, 2013) in tidal creeks in thePlum IslandEstuary,Massachusetts. The chamberswere insertedinto thewater-sediment interface and sealed to limit exchangewith surrounding creekwater (Howard, 2017).Over the ∼4 h daytime incubation period, the concentration of O2 in each chamber increased due to photo-synthetic production exceeding respiration, and the fractionofO2 derived fromphotosynthesis also increased(Figure 2). Based on the initial O2 isotopic composition, observedO2 concentrations in the chambers, and esti-mates of ecosystem production and respiration, it is likely that by the end of the incubation >90% of the O2in each chamber was biological in origin (Howard, 2017), and therefore 17Δ of the dissolved O2 approachedthe biological steady state value (Luz & Barkan, 2000). By the end of each incubation, the value of 17Δ(calculated with 𝜆=0.5179) stabilized at a value near 210 per meg, rather than continuing to increase towardthe value of ∼242 per meg predicted for a steady state system where photosynthetic O2 is derived fromVSMOW (Figure 2).Using an estimated local isotopic composition of H2O based on published data and the salinity in the cham-bers (supporting information Text S1), we predict 17ΔBSS (calculated with 𝜆BSS = 0.5154) to be 215 per megassuming the isotopic fractionation factor for “average phytoplankton” (average of diatoms, cyanobacteria,coccolithophores, and green algae) defined in Luz and Barkan (2011) or 206 per meg assuming a 1:1 mixtureof diatoms and cyanobacteria only (based on microbial cultures from water sampled in the creeks). Thesepredicted values are consistent with the observed 17Δ of ∼210 per meg.If the incubations were extended for a longer duration, all of the atmospheric O2 should eventually beremoved by respiration and the O2 within the chamber should become essentially 100% biological in originand have a composition equivalent to 17ΔBSS (206–215 permeg) (Luz & Barkan, 2011; Nicholson, 2011). If 90%of the O2 in the chamber were biological in origin at the end of the incubation (and the rest derived from gasexchange), the expected 17ΔBSS would be 194–205 per meg.Because the measured 17ΔBSS matched the predicted value with the estimated local isotopic composition ofH2O taken into account, we conclude that chamber incubations are an effective approach for determining17ΔBSS in coastal waters. There are several advantages to the chamber approach compared to published tech-niques for estimating 17ΔBSS (Luz & Barkan, 2011; Luz et al., 1999). For example, the method measures theisotopic composition of dissolvedO2 directly, rather than the composition of gaseous O2 in a headspace equi-librated with water (Luz et al., 1999; Luz & Barkan, 2000), the phytoplankton are grown in oxygenated water(not zero-O2 conditions) (Barkan & Luz, 2011; Eisenstadt et al., 2010; Luz & Barkan, 2011), and we determinethe isotopic composition for an intact microbial community in situ rather than a single autotrophic species inlab-based culture.MANNING ET AL. GOP ESTIMATES INCORPORATING H2O ISOTOPES 4Geophysical Research Letters 10.1002/2017GL074375Figure 2. (a) Observed and predicted measurements of 17Δ during three chamber incubation experiments. The coloredlines show measured values of 17Δ (calculated with 𝜆 = 𝛾R = 0.5179). The gray shaded region shows the predictedvalue of 17ΔBSS (calculated with 𝜆 = 𝜆BSS = 0.5154) using the estimated local H2O isotopic composition, and the dashedblack line shows the 17ΔBSS value for VSMOW water (see text). Error bars indicate uncertainty due to measurementand sampling. (b) Measured O2 concentrations in each chamber. Dash-dotted black line shows the equilibrium O2concentration. Variability in [O2] between chambers is due to expected heterogeneity in benthic estuarine sediments.Notably, 𝛿18O-O2 and 𝛿17O-O2 approached an asymptote in the same manner as17Δ but differed betweenchambers and did not reach the published values for photosynthetic end-members (supporting informa-tion Text S4 and Figure S1). GOP calculations are much more sensitive to 17ΔBSS than to the absolute valueof 𝛿18O-O2P . For example, changing 𝛿18O-O2P from −22‰ to −11‰ changes GOP by less than 5%, as longas 17ΔBSS is held constant. Therefore, the fact that we estimated 17ΔBSS but not 𝛿18O-O2P does not signifi-cantly reduce the utility of the chamber incubation approach and does not affect our conclusion that thelocal TOI composition of H2O should be incorporated intoGOP calculations in systemswhere theH2O isotopiccomposition differs from VSMOW.4. Beaufort Sea, Arctic OceanIn the Arctic Ocean, 𝛿18O-H2O in the surface mixed layer can range from −7 to 0‰ (LeGrande & Schmidt,2006). To quantify the effect of incorporating the local H2O isotopic composition into Arctic Ocean GOP esti-mates, we use a data set of measurements of 17Δ from summer 2011 to 2012 in the Beaufort Sea region ofthe Canada Basin (Stanley et al., 2015) and incorporate measurements of 𝛿18O-H2O from the same cruise. Weassumeafixed 17O-excess of−5permeg, the averagenear-surface seawater value fromLuz andBarkan (2010);however, this value is uncertain because no published observations are available in this region. For this dataset, the calculated GOP assuming the isotopic composition of H2O is VSMOW is on average 29% lower (range48 to 19% lower) than the calculated GOP using the local isotopic composition of H2O (Figures 3 and S2 andTable 1). Increasing or decreasing the 17O-excess by 8 per meg for all samples (two standard deviations ofthe global seawater data set from Luz and Barkan (2010)) changes the calculated GOP by 5–6% on average.The systematic error in Beaufort Sea GOP calculations associated with assuming the H2O isotopic composi-tion is equivalent to VSMOW is larger in magnitude than most other sources of error in the calculation, suchas the gas transfer velocity (10–15% error) (Ho et al., 2006). Given that the isotopic composition of H2O dif-fers from VSMOW throughoutmost of the Arctic Ocean, we expect that GOP estimates in other regions of theArctic Ocean would have similarly large systematic errors if the isotopic composition of H2O is not incorpo-rated into the GOP calculation.For future studies of GOP from 17Δmeasurements in the Arctic Ocean, we strongly recommend investigatorsalsomeasure 𝛿18O-H2O and preferably the17O-excess and that they include these variables in the GOP calcu-lation. Recently, in many regions of the Arctic Ocean the freshwater content has substantially increased andsea ice cover has substantially decreased, and therefore, it would likely be inaccurate to infer the TOI com-position of H2O using previously published, location-based measurements of 𝛿18O-H2O (McPhee et al., 2009;Morison et al., 2012; Yamamoto-Kawai et al., 2009). Additionally, in the Arctic Ocean, there are three salin-ity and H2O isotope end-members (seawater, meteoric water, and sea ice melt) (Eicken et al., 2002; Ekwurzelet al., 2001; Östlund & Hut, 1984; Yamamoto-Kawai et al., 2008). The 17O-excess of these three end-membersMANNING ET AL. GOP ESTIMATES INCORPORATING H2O ISOTOPES 5Geophysical Research Letters 10.1002/2017GL074375-8 -6 -4 -2 0 2δ18O-H2O [per mil vs. VSMOW-SLAP]-60-50-40-30-20-10010% error in GOPVSMOW0 50 100 15017 [per meg vs. air O2]Beaufort SeaBering SeaNorth PacificBellingshausenCaliforniaSouth AtlanticWaquoit Bay-70-60-60-50-40-40-30-30-20-20-10-1000101020-6 -5 -4 -3 -2 -1 0 1 2δ18O-H2O [per mil vs. VSMOW-SLAP]02040608010012014017 [per meg vs. air O2]Beaufort SeaBering SeaNorth PacificBellingshausenCaliforniaSouth Atlantic-60-40-20020% error inGOPVSMOWa           bc-50Figure 3. Error in GOP calculations caused by assuming the water isotopic composition is equivalent to VSMOW(GOPVSMOW) compared to using the local H2O isotopic composition. Percent error as a function of (a) 𝛿18O-H2O and(b) 17Δ for all data sets shown in Table 1. Black lines indicate one standard deviation of the values for each region.(c) Contour plot showing % error in GOPVSMOW as a function of17Δ and 𝛿18O-H2O, for a fixed 17O-excess of −5 per meg(the average oceanic value), calculated using equation (7). Individual samples from each oceanic data set are shown.Table 1Globally Distributed Triple Oxygen Isotope Data and Errors in Calculated GOP Caused by Assuming the H2O Isotopic Composition is Equivalent to VSMOW𝛿18O-H2Oa 17O-excessa(‰ versus (per meg versus 17ΔLocation VSMOW-SLAP) VSMOW-SLAP) (per meg versus air O2) % Error in GOPb ReferenceBeaufort Sea −3.6 (−6.1, −2.3) −5 22 (6, 80) −29 (−48, −19) Stanley et al. (2015)Bering Sea −1.1 (−1.7, −0.7) −5 45 (8, 119) −11 (−16, −7) Prokopenko et al. (2011)California Coast −0.5 (−0.5, −0.3) −5 64 (14, 170) −6 (−11, −5) Munro et al. (2013)Northeast Pacific −0.4 (−0.8, 0.4) −5 34 (11, 67) −4 (−9, 1) Juranek et al. (2012)South Atlantic 0.8 (0.3, 1.4) −5 32 (17, 50) 5 (1, 6) Howard (2017)Bellingshausen Sea −0.5 (−0.7, −0.4) −5 35 (9, 109) −6 (−10, −5) Castro-Morales et al. (2013)Bras d’Or Lake −7.1 (−9.3, −4.8) 16 (6, 27) 35 (9, 54) −46 (−56, −33) Manning (2017)Waquoit Bay −2.9 (−6.1, −1.1) 1 (−5, 11) 100 (45, 236) −30 (−234, −11) This manuscriptNotes. Values are reported as mean (min, max). All tabulated data include only near-surface samples (depth <10 m). aDetails on the estimation of 𝛿18O-H2O and17O-excess for eachdata set are provided in the supporting information Text S1. Analyticalmethods are described in supporting information Text S5 and S6 (Barkanand Luz, 2003; Epstein and Mayeda, 1953; Stanley et al., 2010, 2015). bPercent error in GOP is calculated as % error = (GOPVSMOW − GOP)∕GOP × 100% whereGOPVSMOW is calculated assuming the H2O is equivalent to VSMOW and GOP is calculated using the estimated local H2O isotopic composition.MANNING ET AL. GOP ESTIMATES INCORPORATING H2O ISOTOPES 6Geophysical Research Letters 10.1002/2017GL074375is currently unknown in this region. Therefore, it is not possible to use the two-end-member approach andestimate 𝛿18O-H2O and 𝛿17O-H2O based on salinity alone, nor to estimate 𝛿17O-H2O based on measured𝛿18O-H2O and salinity (supporting information Text S1). If the17O-excess of the three end-members wereknown, one could estimate the 17O-excess using a water mass analysis based on measurements of 𝛿18O-H2Oand an additional freshwater source tracer such as total alkalinity (Yamamoto-Kawai et al., 2008). Concurrentmeasurements of both 𝛿18O-H2O and17O-excess would avoid the uncertainties associated with the watermass analysis.5. Implications for Oceanic and Estuarine SystemsWe use a variety of TOI measurements in other oceanic regions and brackish waters to estimate the error inGOP estimates caused by assuming the water is equivalent to VSMOW (Table 1). In all data sets, the magni-tude of the error in GOP estimates is strongly dependent on the magnitude of 𝛿18O-H2O and more weaklydependent on the magnitude of 17Δ (Figure 3c). In the Bering Sea, GOP is up to 16% lower when calculatedusing VSMOW H2O instead of the local H2O isotopic composition, and in other oceanic regions outside theArctic, the maximum error varies from 11% too low to 6% too high.In freshwater and brackish regions, the magnitude of the errors in calculated GOP is significantly larger, up to56% in an estuary inNova Scotia, Canada, andover 200% in an estuary inMassachusetts (Table 1). These largererrors occur primarily due to the large influence of meteoric water in estuaries, which has a 𝛿18O-H2O of −6to −9‰ in these regions, and in the Massachusetts estuary the errors are exacerbated by high values of 17Δin some samples (Figure 3b). Notably, there are somewhat large uncertainties in the appropriate 17O-excessto use for brackish systems because published data are limited and the range of observed values is large(supporting information Text S2). In future studies, we recommend that authors applying the TOI approachto estimate GOP in regions influenced by meteoric water should measure the 17O-excess of H2O in additionto 17Δ of O2.6. Conclusions and Future OutlookWe have shown that the local TOI composition of H2O is an important variable to include in GOP calculationsin systems where the isotopic composition of H2O differs from VSMOW. We have presented a quantita-tive approach for estimating the local TOI composition of H2O and incorporating this variable into GOPcalculations. Chamber incubation data in an estuary validated this approach and suggest that the TOIapproach could be used to estimate GOP in estuarine and freshwater systems, which could greatly expand itsutility (Staehr et al., 2012).New analytical methods could enable the routine measurement of 𝛿18O-H2O and17O-excess as a part offuture TOI-based GOP studies and promise to increase scientific understanding of spatial trends in H2Oisotopes. Until the past couple of years, analytical methods with sufficient precision to resolve the 17O-excessof H2O were mass spectrometer-based, required a specialized front end and the generation of hazardouschemicals, and did not permit the simultaneous measurement of 𝛿17O and 𝛿2H (Luz & Barkan, 2005).The recent development of commercially available laser-based sensors that simultaneously measure the17O-excess and 𝛿2H of H2O directly, without complicated sample processing (Berman et al., 2013; Steig et al.,2014), will greatly increase the number of laboratories making this measurement and reduce analytical costs(Schauer et al., 2016).Openly accessible databases of 17O-excess in natural waters, similar to existing databases for 𝛿18O-H2O and𝛿2H-H2O (Schmidt et al., 1999; IAEA/WMO, 2016), would be highly beneficial to the stable isotope geochem-istry community. For example, these data could improve estimates of the 17O-excess for data sets where TOImeasurements of O2 and H2O were not collected simultaneously, including previously published data sets.Additionally, public archiving of all data used in TOI-based GOP studies will ensure that GOP can be recalcu-lated in the future based on new consensus about the most appropriate way to perform these calculations(supporting information Text S2).ReferencesAngert, A., Rachmilevitch, S., Barkan, E., & Luz, B. (2003). Effects of photorespiration, the cytochrome pathway, and the alternative pathwayon the triple isotopic composition of atmospheric O2. Global Biogeochemical Cycles, 17(1), 1030. https://doi.org/10.1029/2002GB001933Barkan, E., & Luz, B. (2003). High-precision measurements of 17O/16O and 18O/16O of O2 and O2/Ar ratio in air. Rapid Communications inMass Spectrometry, 17(24), 2809–2814.AcknowledgmentsWe acknowledge funding fromNSF OCE-1029676, OCE-1233678,PLR-1304406, and PLR-1547011 toR.H.R. Stanley, scholarships fromNSERC and CMOS to C.C. Manning,and an NDSEG fellowship toE. M. Howard. We thank MichiyoYamamoto-Kawai for providing theCanada Basin 𝛿18O-H2O data andAmanda Spivak for assisting withincubation chamber design. We thankBoaz Luz and Laurence Yeung for theirthoughtful reviews that improved themanuscript. All data sets collectedby the coauthors that were used inthis study are freely available online:https://doi.org/10.18739/A2W389for Beaufort Sea, https://doi.org/10.5281/zenodo.351764 for WaquoitBay, https://doi.org/10.5281/zenodo.352786 for Bras d’Or Lake, andhttps://doi.org/10.1575/1912/8654 forPlum Island Estuary and the SouthAtlantic. Software for calculatingGOP is available on GitHub(Manning and Howard (2017),http://github.com/caramanning/calcGOP).MANNING ET AL. 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D., Yeung, L. Y., & Kohl, I. E. (2014). On theΔ17O budget of atmospheric O2. Geochimica et Cosmochimica Acta, 135, 102–125.MANNING ET AL. GOP ESTIMATES INCORPORATING H2O ISOTOPES 9 Supporting Information for“Revising estimates of aquatic gross oxygen production by the triple oxygenisotope method to incorporate the local isotopic composition of water”Cara C. Manning1,2,3,4, Evan M. Howard1,2,3,5, David P. Nicholson2, Brenda Y. Ji6, Zoe O.Sandwith2, and Rachel H.R. Stanley61MIT/WHOI Joint Program in Oceanography/Applied Ocean Science and Engineering, Woods Hole, MA, USA.2Department of Marine Chemistry and Geochemistry, Woods Hole Oceanographic Institution, Woods Hole, MA, USA.3Department of Earth, Atmospheric and Planetary Sciences, Massachusetts Institute of Technology, Cambridge, MA, USA.4Department of Earth, Ocean and Atmospheric Sciences, University of British Columbia, Vancouver, BC, Canada.5School of Oceanography, University of Washington, Seattle, WA, USA.6Department of Chemistry, Wellesley College, Wellesley, MA, USA.Code availabilityThe code for calculating gross oxygen production from triple oxygen isotope data ispublished as Manning and Howard [2017]:Manning, C. C., & Howard, E. M. (2017). calcGOP: MATLAB functions for the cal-culation of gross oxygen production from triple oxygen isotope data used in Manning et al.(2017) Geophysical Research Letters, release 1.0. http://github.com/caramanning/calcGOP. https://doi.org/10.5281/zenodo.376786.Contents1. Text S1 to S62. Figures S1 to S2Text S1. Method of estimating the local H2O isotopic compositionFor oceanic systems, we estimate δ 18O-H2O using a published database [Schmidtet al., 1999; LeGrande and Schmidt, 2006], and assume a 17O-excess of -5 (standard devi-ation 4) per meg, the average value for near-surface seawater (<5 m depth) reported in Table1 of [Luz and Barkan, 2010]. For freshwater systems (precipitation-derived or meteoric wa-ter), we use published δ 18O-H2O data for the study region, and a17O-excess of 27 (standarddeviation 20) per meg, or a value specific to the study region where data is available [Luzand Barkan, 2010; Li et al., 2015]. The value of 27 per meg is the average of all publishedfreshwater measurements compiled in Li et al. [2015], Tables 3 and S2.For brackish systems (e.g., estuaries), we assume that the water represents a mixtureof local seawater and local meteoric water. A common approach to evaluate mixing betweendifferent water types in estuaries is to plot salinity versus δ 18O-H2O. Both tracers tend tobehave conservatively and mix as a function of the two endmembers: seawater and meteoricwater [Macdonald et al., 1995; Surge and Lohmann, 2002; Wankel et al., 2006]. We defineδ 18O-H2O and salinity for the two endmembers and then estimate δ18O-H2O for each watersample as a linear function of the measured salinity. Specifically, given that the meteoricwater endmember has salinity (S) = 0, thenδ 18O-H2Osamp =(δ 18O-H2Osw−δ 18O-H2Omw)SswSsamp+δ 18O-H2Omw, (1)Corresponding author: Cara C. Manning, cmanning@whoi.edu–1–where the subscripts sw, mw, and samp refer to seawater, meteoric water, and the sample,respectively. The value of δ 17O-H2O for each endmember is estimated from δ18O-H2O andthe 17O-excess, and then we use equation 1 with δ 17O-H2O to estimate δ17O-H2O for eachsample. We use δ 17O-H2O in the mixing calculation because δ17O-H2O behaves approxi-mately conservatively during mixing and 17O-excess does not.The method of estimating the H2O triple isotopic composition of each dataset in Table1 is described below.Plum Island Estuary (Section 3): For the chamber experiments, we used a freshwa-ter endmember with S = 0, δ 18O-H2O = -6.657h, and 17O-excess = 13.5 per meg [Li et al.,2015] and a seawater endmember with S = 31.5 PSS, δ 18O-H2O = -1h, and 17O-excess =-5 per meg [Schmidt et al., 1999; LeGrande and Schmidt, 2006; Luz and Barkan, 2010]. Weused the observed salinity in the chambers (24.9–28.6 PSS) and equation 1 to estimate thevalue of 17∆BSS for each sample. The local meteoric water endmember δ 18O-H2O and17O-excess is based on measurements of tap water collected in Chicopee, MA and Lewiston, MEfrom Li et al. [2015]. The seawater endmember δ 18O-H2O and salinity are based on mea-surements in nearby North Atlantic surface waters from Schmidt et al. [1999], and the 17O-excess for seawater is the average value for surface waters from Luz and Barkan [2010]Beaufort Sea (Section 4): Values of δ 18O-H2O were measured and the17O-excesswas set to -5 per meg for all samples, the average value for seawater from Luz and Barkan[2010]. As discussed in the main text, there is some uncertainty about the appropriate 17O-excess to use in the Arctic Ocean because there are no published measurements in this regionand because Arctic Ocean seawater is a mixture of three endmembers: ocean water, sea icemelt, and river/meteoric water.Other ocean regions (Section 5, Bering Sea, California Coast, Northeast Pacific,South Atlantic and Bellingshausen Sea): Values of δ 18O-H2O were estimated based onaverage values for the region from a published database and model [Schmidt et al., 1999;LeGrande and Schmidt, 2006], and the 17O-excess was set to -5 per meg for all samples [Luzand Barkan, 2010].Bras d’Or Lake (Section 5): The values of δ 18O-H2O and δ17O-H2O were estimatedusing meteoric water and seawater endmembers with salinity of 0 and 31.25 PSS, δ 18O-H2Oof -9.3 and -1.68h, and 17O-excess of 27 and -5 per meg, respectively [Schmidt et al., 1999;IAEA/WMO, 2016; Li et al., 2015]. The local meteoric water isotopic composition is basedon a multi-year precipitation record from Truro, Nova Scotia (for δ 18O-H2O) and the aver-age 17O-excess value for freshwater given by Li et al. [2015].Waquoit Bay (Section 5): The values of δ 18O-H2O and δ17O-H2O were estimated us-ing meteoric water and seawater endmembers with salinity of 0 and 31.5 PSS, δ 18O-H2O of-6.657 and -1h, and 17O-excess of 13.5 and -5 per meg, respectively [Schmidt et al., 1999;Luz and Barkan, 2010; Li et al., 2015]. The local meteoric water endmembers are basedon measurements of tap water collected in Chicopee, MA and Lewiston, ME from Li et al.[2015].Text S2. Estimating the 17O-excess in other systems and incorporating it into theGOP calculationThe 17O-excess in near-surface oceanic waters varies over a relatively small range,from -15–1 per meg, according to the only published dataset we are aware of, Luz and Barkan[2010]. However, data coverage is sparse and no data has been published for the Arctic Ocean,which has a significant freshwater influence.The 17O-excess of published freshwater measurements varies over a wider range, from-35 to 76 per meg [Li et al., 2015]. For δ 18O-H2O, there are multi-year records of precipi-–2–tation in many regions, whereas published 17O-excess measurements are almost exclusivelysingle time point measurements [Li et al., 2015; IAEA/WMO, 2016]. Given the limited 17O-excess data currently available, we recommend that investigators measure the 17O-excess infuture studies whenever possible.We concur with other authors that in TOI-based GOP studies, all data used to calculateproductivity (e.g., δ 17O-O2, δ18O-O2,17∆, [O2], temperature, salinity, and depth, as wellas the wind speed record used and/or gas transfer velocity, and information on the H2O iso-topic composition if applicable) should be archived and openly available, so that productivitycan be recalculated in the future based on new consensus and information about the most ac-curate equations and values to use in these calculations [Hamme et al., 2012; Juranek andQuay, 2013].Text S3. Choice of λ in GOP calculationsWe contend that it is most appropriate to use λ = γR = 0.5179 rather than λBSS whencalculating the rate of gross photosynthetic production in equation 7 [Kaiser, 2011; Nichol-son, 2011; Nicholson et al., 2012; Hamme et al., 2012; Juranek and Quay, 2013]. However,we acknowledge that some other works have advocated for the use of λBSS in GOP calcula-tions [Luz and Barkan, 2011]. If λBSS is used in equation 7, the calculated GOP is even moresensitive to the H2O isotopic composition than presented in this work.Text S4. Behaviour of δ 18O-O2 and δ 17O-O2 in chamber incubationsIn the chamber experiments, δ 18O-O2 and δ17O-O2 asymptote but do not reach thepublished photosynthetic endmember values, which are based on laboratory cultures of indi-vidual autotrophic species (Figure S1). For example, δ 18O-O2 at the end of each incubationranged from -11 to -15.5h in the three chambers shown here, whereas the predicted valuefor photosynthetic O2 derived from ‘average phytoplankton’ with the estimated isotopiccomposition of H2O was ∼-22h. We suspect that these differences occurred because weincubated a heterogeneous, intact microbial community (both autotrophs and heterotrophs),and also because of differences in biogeochemical conditions between chambers, and theshort incubation period. The heterogeneous microbial community would affect the finalvalue of δ 18O-O2 and δ17O-O2 but not17∆BSS because there are a wide range of respira-tory fractionation factors (18εR) reported in the literature, whereas the value of γR (the ratioof 17εR/18εR) is nearly universal for a wide range of organisms [Luz and Barkan, 2005].Text S5. Method for O2 isotope analysisSamples for analysis of the triple oxygen isotopic composition of O2 were collectedfrom Plum Island Estuary (chamber incubations), the Canada Basin, South Atlantic Ocean,Bras d’Or Lake, and Waquoit Bay and analyzed at Woods Hole Oceanographic Institutionwith a Thermo Finnigan MAT 253 isotope ratio mass spectrometer. The analytical proce-dure followed published methods [Barkan and Luz, 2003] with modifications as described inStanley et al. [2010, 2015]. All measurements are referenced to O2 in air collected locally inWoods Hole, MA. Standards of air-equilibrated distilled water are run daily to check instru-ment performance and precision. For all data sets the value of 17∆ for these air-equilibratedsamples was very close to 8 per meg, which is equivalent to the value reported in Stanleyet al. [2010]. The standard deviation of replicate field samples as well as replicate equili-brated water standards was ∼5 per meg for all data sets.Other data sets shown in Table 1 and Figure 2 were analyzed by similar methods [Prokopenkoet al., 2011; Juranek et al., 2012; Castro-Morales et al., 2013; Munro et al., 2013].–3–Text S6. Method for H2O isotope analysisSamples for analysis of δ 18O-H2O were collected from the Arctic Ocean and analyzedat the Oregon State University College of Earth, Ocean and Atmospheric Sciences StableIsotope Laboratory with a Thermo Finnigan DeltaPlus XL isotope ratio mass spectrometer,using the H2O/CO2 equilibration method [Epstein and Mayeda, 1953]. Replicate analyseshad a precision of 0.05h or better.Figure S1.08 10 12 14 16Hour of day-20-15-10-50δ17 O-O2 [‰ vs. air O 2]Chamber 1Chamber 2Chamber 308 10 12 14 16Hour of day-20-15-10-50δ18 O-O2 [‰ vs. air O 2]a   bCaption for Figure S1: Chamber incubation measurements of (a) δ 17O-H2O, and (b)δ 18O-H2O. The samples are from the same dataset plotted in Figure 1. The dashed line oneach figure shows the predicted isotopic composition of O2 produced by “average phyto-plankton" as reported in [Luz and Barkan, 2011], based on culture experiments, and derivedfrom H2O with the isotopic composition estimated as described in the main text.–4–Figure S2.Caption for Figure S2: Comparison of GOP estimates for the Canada Basin based onassuming the water isotopic composition is VSMOW (x-axis) and using the measured δ 18O-H2O and an assumed17O-excess of 0 per meg versus VSMOW-SLAP (y-axis). The blackdashed line is a 1:1 line and all measurements fall above this line. The gas transfer velocityis calculated based on wind speed and as a function of sea ice cover as described in Stanleyet al. [2015]. The blue circles and red diamonds represent data collected in summer 2011and 2012, respectively.ReferencesBarkan, E., and B. Luz (2003), High-precision measurements of 17O/16O and 18O/16O of O2and O2/Ar ratio in air, Rapid Commun. 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